Alfred Wegener’s concept of continental drift, reformulated in the modern theory of plate tectonics, arose in part as a way to explain the geographic distribution of paleoclimate indicators in ancient rocks. Permo-Carboniferous (~300-million-year-old) glacial deposits in distinctly nonpolar regions of present-day Africa, South America, and Australia rectify to polar latitudes when the ancient supercontinent of Gondwana is reconstructed. Continental drift transported them through broad climate belts—humid tropics, arid subtropics, moist and cool temperate zones, and cold and arid polar regions (Box 2.1). Paleoclimate reconstructions, however, reveal that although paleogeography and the plate tectonics that control continental configurations are important, they are not the major determinant of climate change. Global warm climates have prevailed when large continents covered the poles, and deep “snowball Earth” glaciations occurred when there apparently were no polar continents. Instead, it appears that the greenhouse gas content of the atmosphere was the key factor in determining whether a particular interval of Earth’s past was an icehouse or a greenhouse.

Although most deep-time greenhouse climates occurred when there were distinctly different continental configurations, and thus are not direct analogues for the future, past warm climates and abrupt transitions into even hotter states (known as hyperthermal events; Thomas et al., 2000) provide important insights into how physical, biogeochemical, and biological processes operate under warm conditions more analogous to what is anticipated for the future than the moderate and stable climates of the Holocene (past 10,000 years) or the relatively warm interglacials of

Citation Manager

Below are the first 10 and last 10 pages of uncorrected machine-read text (when available) of this chapter, followed by the top 30 algorithmically extracted key phrases from the chapter as a whole.Intended to provide our own search engines and external engines with highly rich, chapter-representative searchable text on the opening pages of each chapter.
Because it is UNCORRECTED material, please consider the following text as a useful but insufficient proxy for the authoritative book pages.

Do not use for reproduction, copying, pasting, or reading; exclusively for search engines.

OCR for page 26
2
Lessons from Past Warm Worlds
Alfred Wegener’s concept of continental drift, reformulated in the
modern theory of plate tectonics, arose in part as a way to explain the
geographic distribution of paleoclimate indicators in ancient rocks.
Permo-Carboniferous (~300-million-year-old) glacial deposits in distinctly
nonpolar regions of present-day Africa, South America, and Australia
rectify to polar latitudes when the ancient supercontinent of Gondwana is
reconstructed. Continental drift transported them through broad climate
belts—humid tropics, arid subtropics, moist and cool temperate zones,
and cold and arid polar regions (Box 2.1). Paleoclimate reconstructions,
however, reveal that although paleogeography and the plate tectonics that
control continental configurations are important, they are not the major
determinant of climate change. Global warm climates have prevailed
when large continents covered the poles, and deep “snowball Earth” gla -
ciations occurred when there apparently were no polar continents. Instead,
it appears that the greenhouse gas content of the atmosphere was the key
factor in determining whether a particular interval of Earth’s past was an
icehouse or a greenhouse.
Although most deep-time greenhouse climates occurred when there
were distinctly different continental configurations, and thus are not
direct analogues for the future, past warm climates and abrupt transi-
tions into even hotter states (known as hyperthermal events; Thomas et
al., 2000) provide important insights into how physical, biogeochemical,
and biological processes operate under warm conditions more analogous
to what is anticipated for the future than the moderate and stable climates
of the Holocene (past 10,000 years) or the relatively warm interglacials of
26

OCR for page 26
27
LESSONS FROM PAST WARM WORLDS
BOX 2.1
Continental Drift and Climate
Plate tectonics has been rearranging Earth’s configuration of continents
ever since the plates on Earth became rigid approximately 2.5 billion years
ago (Figure 2.1). On long (millions of years) timescales, the movement of
tectonic plates—and the continents that ride upon them—has strongly
influenced Earth’s distribution of solar insolation, ocean and atmospheric
circulation, and carbon cycling between the Earth’s deep and shallow
reservoirs, thereby profoundly impacting global climate, sea level, and the
overall planetary ecology.
The arrangement of the continents through time is most reliable for the
past 800 million years, the period for which the chronostratigraphic tools
necessary for reconstructions are available. The global views presented in
Figure 2.1 show how the continents on Earth’s surface may have appeared
during four intervals of time that are noted throughout this report: the unipo-
lar glaciated Pennsylvanian (300 million years ago [Ma]), mid-Cretaceous
(105 Ma), Eocene (50 Ma), and mid-Pliocene (3 Ma).
The major transitions between climatic icehouse and greenhouse con-
ditions are ultimately most probably driven by the deep Earth processes
of plate tectonics, as a function of the long-term balance between CO2
degassing at spreading centers and the conversion of atmospheric CO2 to
mineral carbon through long-term silicate weathering and oceanic carbon-
ate formation (Berner, 2004). For example, the eruptions of large igneous
provinces in the mid-Cretaceous and the subduction of the carbonate-rich
tropical Tethys Sea in the early Cenozoic are the most likely cause of the
high-CO2 equilibrium climates of the Cretaceous and Eocene greenhouses.
Conversely, uplift of the Himalayas and Tibetan Plateau associated with
“docking” of the Indian subcontinent with Asia (~40 Ma), and the evolu-
tion of vascular land plants in the early Paleozoic (~450 Ma), led to the
sequestration of atmospheric CO2 through enhanced weathering of silicate
minerals (Ruddiman, 2007; Archer, 2009).
continued

OCR for page 26
28
BOX 2.1 Continued
FIGURE 2.1 Continental configurations for the Pennsylvanian (upper left), mid-Cretaceous (upper right),
Eocene (lower left), and mid-Pliocene (lower right). Topography was defined on the basis of digital elevation
maps of modern Earth from the U.S. Geological Survey; colors portray climate and vegetation distribution
based on a synthesis of all geological literature relevant to each time slice.
SOURCE: Courtesy R.C. Blakey, Colorado Plateau Geosystems.
UNDERSTANDING EARTH’S DEEP PAST

OCR for page 26
29
LESSONS FROM PAST WARM WORLDS
the Pleistocene (past 2 million years). The following sections describe the
insights provided by understanding past warm periods, including the
role of greenhouse gases in controlling—or “forcing”—global warming;
the impact of warming on ice sheet stability, sea level, and oceanic and
hydrological processes; and the consequences of global warming for eco -
systems and the global biosphere.
CLIMATE SENSITIVITY TO INCREASING CO2
IN A WARMER WORLD
Fundamentally, Earth’s climate results from the balance between
absorbed energy from the sun and radiant energy emitted from Earth’s
surface, with changes to either component resulting in a forcing of the
climate system. The net forcing of the climate system over geological
time caused episodes of warming and cooling that are coincident with
greenhouse and icehouse climates, respectively. Most projections indicate
that, by the end of this century, climate forcing resulting from increased
CO2 will be at least of the same magnitude as that experienced in the
early Cenozoic (during the late Eocene, ~34 Ma) (Figure 2.1), and possibly
analogous to estimates for the Cretaceous Period (~80-120 Ma)—probably
one of the times of greatest radiative forcing since the evolution of animals
(Hay, 2010).
Climate sensitivity—the equilibrium warming resulting from a dou-
bling of atmospheric carbon dioxide relative to preindustrial levels of
CO2—provides a measure of how the climate system responds to external
forcing factors and is also used to compare global climate model outputs
to understand why different models respond to the same external forcings
with different outputs. Climate sensitivity to CO2 strongly influences the
magnitude of warming that Earth will experience at any particular time
in the future (Box 2.2). The magnitude of climate sensitivity and Earth’s
surface temperature are determined by a myriad of short-term (human
timescales) and long-term (thousands to tens of thousands) interactions
and feedbacks (e.g., water vapor, cloud properties, sea ice albedo, snow
albedo, ice sheet and terrestrial biome distribution, ocean-atmosphere CO2
interaction, and silicate weathering).
As noted above, synthesis of the various estimates of Earth’s climate
sensitivity for the past 20,000 years has lead to the general conclusion that
sensitivity most probably lies in the range of 1.5 to 4.5°C (IPCC, 2007), with
some recent projections suggesting that the value may be even as high as
6-8°C (Hansen et al., 2008; Knutti and Hegerl, 2008). However, estimates of
equilibrium climate sensitivity averaged over tens to hundreds of millen -
nia (i.e., long term) and extending back for 400 million years are minimally
between 3 and 6°C (Royer et al., 2007). For the most recent period of global

OCR for page 26
30 UNDERSTANDING EARTH’S DEEP PAST
FIGURE 2.2 Estimated atmospheric pCO2 for the past 45 million years (late Eocene
through Miocene) calculated using all available stable carbon isotopic values of
diunsaturated alkenones in deep-sea sediments. Values of CO2aq were translated to
atmospheric pCO2 using Henry’s Law and a range of dissolved phosphate values
and sea surface temperatures for each site, and a salinity of 35 parts per thousand.
The dark gray shaded region shows the range of maximum to intermediate esti-
mates, and the dashed line represents minimum estimates. The uncertainty in
pCO2 estimates ranges from ~20 percent for the Miocene to 30 to 40 percent for the
Paleogene. The broad pale red band (pCO2 values of 600-1,100 parts per million
by volume) encompasses most of the CO2 concentration range for nonmitigation
emission scenarios projected for the end of this century (figure 10.26 in IPCC, 2007);
the dark red band (values of 800-1,000 ppmv) corresponds to the Intergovernmen -
tal Panel on Climate Change (IPCC) A2 “business-as-usual” scenario.
SOURCE: Modified after Pagani et al. (2005).
warming, the middle Pliocene (~3.0-3.3 Ma), climate sensitivity may have
been as high as 7-9.6°C ± 1.4°C per CO2 doubling (Pagani et al., 2010).
Such values, which are well above short-term climate sensitivity estimates
based on more recent paleoclimate and instrumental records, indicate that
long-term feedbacks operating at accelerated timescales (decadal to cen -
tennial) promoted by global warming can substantially magnify an initial
temperature increase.
As Earth moves toward a warmer climate state, it is important to under-
stand the extent to which climate sensitivity will change due to processes

OCR for page 26
31
LESSONS FROM PAST WARM WORLDS
BOX 2.2
Why Does Climate Sensitivity Matter?
For any particular increase in atmospheric CO2 (and other greenhouse
gases), a system with high climate sensitivity to CO2 will warm more in the
future than a world with low climate sensitivity. Thus, if the climate sensitiv-
ity is high, restricting future global warming will require a larger reduction
in future CO2 emissions than if climate sensitivity is lower. Comparison
of emission scenarios for the period until 2100, calculated for a range of
CO2 stabilization targets (Figure 2.3A) and the corresponding equilibrium
global average temperature increases (Figure 2.3B; IPCC, 2007), based on
the Intergovernmental Panel on Climate Change (IPCC) range of climate
sensitivities (2 to 4.5°C), illustrates the impact of fossil carbon emissions on
future surface temperatures and the extent of reductions required to limit
the warming to ≤2°C relative to preindustrial conditions.
Even if anthropogenic carbon emissions to the atmosphere are reduced,
CO2 levels will continue to increase for a century or more because the
removal of CO2 from the atmosphere by natural processes of carbon seques-
tration (e.g., CO2 absorption by the surface ocean, CO2 fertilization of ter-
restrial vegetation) is slow (Archer et al., 2009). Consequently, temperature
increases may continue for several centuries until equilibrium temperatures
are reached, especially for higher CO2 stabilization targets. However, even
if climate sensitivity is at the lower end of the possible range, global tem-
perature increases of ≥2°C will be reached with CO2 stabilization levels of
450-550 parts per million by volume. Given that equilibrium temperature
increases may be protracted, emissions could continue to increase into
the middle of this century (Figure 2.3; Caldeira et al., 2003). However, if
the climate sensitivity is 4.5°C or greater, then a significant and immediate
reduction in CO2 emissions—to levels ultimately below those of the present
day—is required to stay below a target warming of 2°C.
continued

OCR for page 26
32 UNDERSTANDING EARTH’S DEEP PAST
BOX 2.2 Continued
World CO2 emissions (GtCO2/yr)
A Year
FIGURE 2.3 Global CO2 emissions and equilibrium global average
temperature increases for a range of target CO2 stabilization levels.
(A) Measured (1940 to 2000) and projected (colored shading; 10th to 90th
percentile) global CO2 emissions for the range of IPCC emission scenarios
and associated stabilization CO2 levels indicated by roman numerals
(ppm CO2-eq). (B) Corresponding relationship between the different
that have not operated in recent icehouse climate regimes. One simple
example is to consider a warm world with no sea ice at either pole—as CO2
increases, the sea ice albedo feedback is removed and therefore this nega-
tive feedback’s contribution to climate sensitivity is absent. In addition, it
is important to determine the potential for nonlinear responses that are
specific to a greenhouse or transitional world, and whether such responses
would enhance climate sensitivity. For example, the destabilization of con-
tinental ice sheets resulting from warming of polar regions can potentially
lead to a decrease in deep-water formation, thereby affecting global ocean
circulation, stratification, and carbon cycling, leading to higher climate sen-
sitivity than indicated by present estimates. In sufficiently warm climates,
even water vapor has a nonlinear dependence on temperature, and this can
introduce new and potentially rapid feedbacks, operating at a subdecadal
scale, into the climate system. Destabilization of methane and its release

OCR for page 26
33
LESSONS FROM PAST WARM WORLDS
E quilibrium global average temperature
increas e above preindustrial (°C)
B G HG concentration s tabilis ation level (ppm C O2 -eq)
CO2 stabilization targets shown in (A) and equilibrium global average
temperature increase above preindustrial levels. Colored regions for
each stabilization target were calculated for a range of climate sensitivity
(2–4.5°C) and “best estimate” climate sensitivity of 3°C (blue solid line
in middle of shaded area).
SOURCE: IPCC (2007, Figure 5.1, page 66).
into the atmosphere in response to warming— through either the melting
of terrestrial permafrost reservoirs or the dissolution of subseafloor clath-
rate deposits—would dramatically increase greenhouse gas contents in the
atmosphere. Hence, an initial warming from greenhouse gases released by
burning fossil fuels could end up releasing even more greenhouse gases
from natural sources, exacerbating the original warming of the atmosphere.
TROPICAL AND POLAR CLIMATE STABILITY AND
LATITUDINAL TEMPERATURE GRADIENTS
IN A WARMER WORLD
With more than half of Earth’s surface lying within 30° latitude of
the equator, the response of tropical climates to increased greenhouse gas
forcing is critically important. Modern observational data (Ramanathan and

OCR for page 26
34 UNDERSTANDING EARTH’S DEEP PAST
Collins, 1991) suggest that western tropical Pacific sea surface temperatures
rarely exceed ~30-32°C, and this has led to speculation that the Earth’s
tropics have a “thermostat” that limits maximum sea surface temperatures.
Explanations about how such a thermostat might work have included the
buildup of clouds that reflect heat back into space (Ramanathan and Collins,
1991), evaporative cooling (Hartmann and Michelsen, 1993; Pierrehumbert,
1995), winds, or an increase in transport of heat out of the tropics by ocean
currents (Clement et al., 1996; Sun and Liu, 1996). Most of these studies have
used present-day data to explain surface temperature regulation, although
there are artifacts in these datasets that call into question the robustness of
the observed trends (Clement et al., 2010). Paleoclimate reconstructions of
tropical temperatures during past greenhouse times, however, document
sea surface temperatures that were much warmer than modern tropical
maxima—possibly as high as 42°C—and thus were probably not thermo-
statically “regulated” (Bice et al., 2006; Came et al., 2007; Pearson et al., 2007;
Trotter et al., 2008; Kozdon et al., 2009).
The discovery of a giant Paleocene snake fossil in South America
(Head et al., 2009; Huber, 2009; although see discussions by Makarieva et
al., 2009; Sniderman, 2009), as well as other terrestrial paleotemperature
indicators such as paleoflora leaf-margin analysis and stable isotope com -
positions of biogenic apatites and soil minerals (Fricke and Wing, 2004;
Tabor and Montañez, 2005; Passey et al., 2010), further suggests anoma-
lously high continental temperatures (~30-34°C) for the terrestrial tropics
of past warmer worlds. Additionally, coupled climate model simulations
with large radiative forcings and/or paleoclimate simulations for elevated
greenhouse gases do not produce a thermostatic regulation of tropical
temperatures (e.g., Boer et al., 2005; Poulsen et al., 2007b; Cherchi et al.,
2008), suggesting that the tropical warming in response to greenhouse gas
forcing is neither moderated nor local in its impacts (Xie et al., 2010). Such
deep-time paleoclimate studies have documented that tropical surface
temperatures during past greenhouse periods were not thermostatically
regulated by the negative feedback processes that operate in the current
icehouse climate system, further illustrating how knowledge of deep-time
warm periods is fundamental to understanding Earth’s climate system.
There is also abundant evidence for anomalous polar warmth during
past greenhouse periods (e.g., middle Cretaceous to Eocene, Pliocene; see
Figure 2.4) associated with reduced equator-to-pole temperature gradients
(e.g., Huber et al., 1995; Crowley and Zachos, 2000; Hay, 2010; Miller et al.,
2010). To date, climate models have not been able to simulate this warmth
without invoking greenhouse gas concentrations that are notably higher
than proxy estimates (Figure 2.4; Bice et al., 2006). This has prompted mod-
eling efforts to explain high-latitude warmth through vegetation (DeConto
et al., 1999), clouds (Sloan and Pollard, 1998; Abbot and Tziperman, 2008;

OCR for page 26
36 UNDERSTANDING EARTH’S DEEP PAST
Kump and Pollard, 2008), intensified heat transport by the oceans (Barron
et al., 1995; Korty et al., 2008), and increased tropical cyclone activity (Sriver
and Huber, 2007; Fedorov et al., 2010). The ability to successfully model a
reduced latitudinal temperature gradient state, including anomalous polar
warmth, presents a first-order check on the efficacy of climate models as
the basis for predicting future greenhouse conditions.
Since significant changes in tropical and polar surface temperatures
and pole-to-equator temperature gradients occurred in the past, and could
occur in a future warmer world, it is imperative to understand the mecha-
nisms and feedbacks that lead to such changes and their consequences for
atmospheric and oceanic circulation (Hay, 2008). The fundamental mis -
match between model outputs, modern observations, and paleoclimate
proxy records discussed above, however, may indicate some very impor-
tant deficiencies in scientific knowledge of climate and the construction of
climate models (e.g., Huber, 2008). Resolution of this disparity, as well as
an improved understanding of the anthropogenic signal in observational
data, can likely be obtained by analysis of paleoclimate records from past
warm worlds.
HYDROLOGICAL PROCESSES AND THE GLOBAL WATER CYCLE
IN A WARMER WORLD
Earth’s hydrological processes—including precipitation, evaporation,
and surface runoff—are susceptible to, and play a critical role in, both
past and future climate change (Pierrehumbert, 2002). Large-scale atmo-
spheric processes determine the general position of climate zones and
the intensity of precipitation and storms; the intertropical convergence
zone is a region of significant rainfall, while large regions of atmospheric
subsidence lead to dry desert regions. Regional hydroclimates, such as
the Southwest Indian and the East Asian summer monsoons, which affect
nearly half of Earth’s human population, are highly sensitive to distal
climate changes and to mean warming (Sinha et al., 2005; Wang et al.,
2005) via teleconnections (e.g., changes in high-latitude surface tempera-
tures or Arctic sea ice extent impact lower-latitude climate through atmo-
spheric processes). Overall, the vapor-holding capacity of the atmosphere
increases substantially with increased global mean temperatures if there
is no change in the relative humidity. Consequently, climate models for
global warming predict an intensified hydrological cycle and, on a global
scale, enhanced precipitation (IPCC, 2007).
Observations over the past few decades indicate that precipitation
has increased faster (~7 percent per degree of surface warming; Wentz et
al., 2007) than that predicted by models (1-3 percent per degree of surface
warming; Zhang et al., 2007). Although the reasons for this substantial

OCR for page 26
52
Multiple hypoxia events—Mediterranean
Thickness (cm)
50 25 0
75
125 100
Ocean Anoxic Event 1B, Western North Atlantic
Thickness (cm)
50 25 0
75
125 100
FIGURE 2.11 Examples of ancient hypoxic episodes in a Plio-Pleistocene drill core from the Mediterranean. Black bands in the
upper photo (from Ocean Drilling Program Site 964) are “sapropels”—layers rich in organic carbon—formed when the surface
waters of the Mediterranean abruptly warmed, became fresher, and ceased to circulate as they do today. Warming and high nutri -
ent supply led to blooms of algae and bacteria, preserved as a layer of organic carbon on the seafloor. Lower photo shows a core
from the western North Atlantic (from ODP Site 1049) in which a similar layer rich in organic carbon was deposited during an
event 112 million years ago that was broadly analogous to the Mediterranean hypoxic events.
SOURCE: Images courtesy Integrated Ocean Drilling Program Science Services.

OCR for page 26
53
LESSONS FROM PAST WARM WORLDS
continue to increase well above the levels of the Pleistocene interglacials
and as the geographic distributions of climate zones change. For example,
higher CO2 levels are expected to saturate the CO2 fertilization effect,
resulting in a shift of the terrestrial biosphere from a net sink to a net
source of carbon sometime within this century (Cao and Woodward, 1998;
Cox et al., 2000). Furthermore, as the surface oceans warm and become less
alkaline with increasing atmospheric CO2, carbonate-bearing animals will
be strongly impacted (e.g., see Box 2.6), further perturbing biota-climate
feedbacks compared with those reconstructed from the recent past.
The deep-time geological record, in particular the record of warm
periods of higher atmospheric pCO2 and including the transitions into
and out of these periods, has the potential to yield unique insights into the
nature and rate of biotic response to climate perturbation as well as into
the biota-climate feedbacks accompanying global warming. For example,
the mid-Paleozoic “greening” of continents, marked by the evolution and
spread of vascular land plants (Gensel and Andrews, 1987; Beerbower et
al., 1992), records a large-scale natural experiment in the climatic effects
of vegetation—reflecting the contrast between a largely unvegetated pre-
Devonian world compared with one that was heavily vegetated—that
has been linked to major changes in atmospheric CO2 and a vastly dif-
ferent hydrological regime (Algeo et al., 1995, 2001). Another example of
the potential of the deep-time record is provided by the repeated major
restructuring and turnover within terrestrial floral communities that
occurred in step with recurrent shifts in surface temperature, precipita-
tion levels, seasonality, and soil moisture during the demise of the Late
Paleozoic Ice Age at ~295-260 Ma, the vegetated Earth’s only analogue of
a CO2-forced icehouse-to-greenhouse transition (see Box 2.7).
More recently, the gradual but extreme warming (perhaps up to >30
to 42°C in the tropics) of the early Eocene greenhouse (Box 2.8) may have
triggered a major tropical vegetation die-off, with substantial changes in
evapotranspiration fluxes, precipitation, albedo, surface temperature, and
carbon feedbacks (Huber, 2008). During the transient global warming and
short-term aridity of the Paleocene-Eocene Thermal Maximum (PETM)
major restructuring among terrestrial biomes resulted in expansion in the
latitudinal range of subtropical and tropical rainforests (Wing et al., 2005).
Oxidation of the terrestrial biosphere at the Paleocene-Eocene boundary
may have released several gigatons of carbon into the atmosphere, sub -
stantially amplifying the existing greenhouse warming and its climate
effects (Kurtz et al., 2003).
The potential vulnerability of modern biotic communities to cata-
strophic disruption (Jackson et al., 2001; Chase and Leibold, 2003) is an
issue designated as one of the “grand challenges” in the environmental
sciences (NRC, 2001). Globally, current extinction rates are estimated to

OCR for page 26
54 UNDERSTANDING EARTH’S DEEP PAST
BOX 2.6
Impact of Past and Future Climate Change on Coral Reefs
Healthy coral reef ecosystems develop under a relatively narrow range
of ocean temperatures and chemistry (Kleypas et al., 1999) and are there-
fore sensitive indicators of environmental conditions. Global change
models predict that reef systems, with their abundant biodiversity, will
be exposed to higher ocean temperatures and increasingly more acidic
waters in the next century (Hoegh-Guldberg et al., 2007; see Figure 2.12).
Indeed, research suggests that global climate change has already caused
steep declines in coral growth on reef systems around the world (Hoegh-
Guldberg, 1999). Culturing experiments with corals in acidified waters
show that skeleton growth drops as acidity increases and, in extreme cases,
coral colonies can lose their skeletons completely and grow as soft-bodied
anemone-like animals (Fine and Tchenov, 2007). In fact, ocean acidification
may vie with global warming as the most severe threat to marine ecosystems
(Hoegh-Guldberg et al., 2007; De’ath et al., 2009). Reef systems, however,
are intrinsically complex structurally and ecologically, making it difficult to
evaluate the likely impact of future global change on modern reefs based
solely on studies of present-day systems. The geologic record of fossil reef
evolution provides opportunities to study the response of reef ecosystems
to past episodes of increased global temperatures and ocean acidification.
The coral reef crisis occurring in modern oceans may be the sixth such
major reef crisis recorded in the past 500 million years of marine metazoan
evolution. Four of the previous five metazoan reef crises appear to have
been driven by greenhouse gas-forced global warming that was probably
associated with ocean acidification (Veron, 2008; Kiessling and Simpson,
2010). At least three of these reef crises were associated with massive re-
lease of greenhouse gases into the oceans and atmosphere, leading to pCO2
increases analogous to—or perhaps even greater than—those anticipated
for Earth’s future. For example, major reef crises during the Early Jurassic
and during the Cretaceous were associated with massive releases of vol-
canic CO2 to the atmosphere that led to global warming, oceanic anoxia,
and quite likely ocean acidification (Knoll et al., 1996; Svensen et al., 2007;
Hermoso et al., 2009). One of the major reef crises occurred at the same
time as the best-documented case of greenhouse gas-induced ocean acidi-
fication in the geological record, the Paleocene-Eocene Thermal Maximum
(PETM) of 56 Ma (described in more detail in the next chapter). Although
coral-algal reefs began to decline throughout the Tethyan region in the
early Eocene due to the development of very warm (~30-35°C) tropical
sea surface temperatures (Scheibner and Speijer, 2008) (Figure 2.13), PETM
extinction rates indicate that ocean acidification must have been a major
continued

OCR for page 26
LESSONS FROM PAST WARM WORLDS
FIGURE 2.12 Extant examples of reefs from the Great Barrier Reef that are used as analogues for the ecological
structures anticipated for atmospheric CO2 values of (A) 380 ppmv, (B) 450-500 ppmv, and (C) >500 ppmv (Hoegh-
Guldberg et al., 2007). Scenario C corresponds to a +2°C increase in sea temperature. The atmospheric CO 2 and
temperature increases shown are those for the scenarios and do not refer to the particular locations photographed.
SOURCE: Photographs by and with permission of Ove Hoegh-Guldberg, Global Change Institute, University of
Queensland.
55

OCR for page 26
57
LESSONS FROM PAST WARM WORLDS
BOX 2.7
Climate-Driven Restructuring of Late Paleozoic Tropical Forests
Integration of climate proxy records with tropical paleobotanical
archives from the Late Paleozoic shows repeated climate-driven ecosystem
restructuring of paleotropical flora in step with climate and pCO2 shifts,
illustrating the biotic impact associated with past CO2-forced turnover
to a permanent ice-free world (Montañez et al., 2007; DiMichele et al.,
2009). Wetland flora—consisting of ferns and pteridosperms, sphenopsids,
and lycopsids—was rapidly replaced in the earliest Permian by dryland
flora that diversified in the now seasonally dry habitats created by an
abrupt shift from ever-wet to semiarid conditions. Tree and fern-rich floras
reappeared during wetter, cooler conditions of the subsequent glaciation
at ~285 Ma characterized by lowered pCO2 (Figure 2.14). Such dramatic
floristic changes occurred with each climate transition during the final stage
of the Late Paleozoic Ice Age.
The fact that these temporally successive floras tracked climatic condi-
tions and contained progressively more evolutionarily advanced lineages
suggests that evolutionary innovation occurred in extrabasinal areas and
was revealed by climate-driven floral migration into lowland basins. One
such scenario occurred during the return to cold conditions at the close of
the Early Permian when unique seed-plant assemblages, not observed again
until the Late Permian (conifers) and Mesozoic (cycads), migrated into low-
land basins. Climate transitions drove macroevolution in the oceans as well,
including significant changes in marine invertebrate biodiversity coincident
with the appearance of a diverse array of early terrestrial vertebrate lineages
and major restructuring of floral biomes (Clapham and James, 2007).
continued

OCR for page 26
58 UNDERSTANDING EARTH’S DEEP PAST
BOX 2.7 Continued

OCR for page 26
59
LESSONS FROM PAST WARM WORLDS
FIGURE 2.14 Floral abundance patterns (A and B) in the paleotropics
during the latest Pennsylvanian through Middle Permian, plotted against
(C) estimated pCO2 (blue line) and paleo-sea surface temperatures (red
band). Periods of glaciation or widespread cooling in the high southern
latitudes are shown by blue bars. Top panel (A) shows the temporal
distribution of typical latest Carboniferous wetland floras (ferns and
pteridosperms, sphenopsids, and lycopsids). The middle panel (B) illus -
trates the temporal distribution of dryland floras (conifers, callipterids
and other seed plants) that diversified in seasonally dry habitats. The
short-term intercalation of the two floras—at the likely millennial scale—
occurred with the return of wetland floras in the mid Early Permian
transient glaciation under cooler and wetter conditions brought on by
significantly lowered pCO2 and renewed glaciation.
SOURCE: Modified after Montañez et al. (2007).

OCR for page 26
60 UNDERSTANDING EARTH’S DEEP PAST
BOX 2.8
Biome Distributions in the Cretaceous-Early Eocene Hothouse
The discovery more than a century ago of coal seams, fossil forests, and
fossil leaves of warm temperate trees above the Arctic Circle on the west
coast of Greenland, more than 1,000 km north of the modern tree line,
was an early indicator of anomalous warmth at high latitudes in the past.
That arctic regions more than 50 Ma had been forested as far north as there
was land, despite a continental configuration similar to that of the present
day, became increasingly more apparent with the discovery of hundreds
of similar sites on arctic islands and across arctic Asia and North America
(Spicer et al., 2008). Northern hemisphere Cretaceous-early Eocene polar
forests were fundamentally different from present-day boreal forests, which
do not grow north of the Arctic Circle, as they are dominated by deciduous
conifers related to the bald cypress and dawn redwood and by a variety of
deciduous broadleaf trees. Leaf margin analysis of Paleocene-Eocene floras
from arctic Canada on Axel Heiberg Island (at 78ºN) yield mean annual
temperatures of 10º ± 2ºC (Basinger et al., 1994), in striking contrast to the
modern mean annual temperatures of minus 30ºC. The fossil floral record
indicates that in this warmer world, both subtropical and tropical rainforests
had greatly expanded latitudinal ranges (Figure 2.15).
Subtropical conditions in the polar Arctic are further indicated by the
occurrence of early Eocene crocodiles, turtles, and snakes on Ellesmere
Island at 80ºN at this time (Dawson et al., 1976; Markwick, 2007). Sub-
sequent discovery of fossil mammals and plants, related to contemporary
biotas in France and Wyoming, confirmed the hypothesis that Arctic Canada
at this time was part of a warm temperate land connection between Europe
and North America (Hickey et al., 1983). The recent and surprising dis-
covery of the aquatic fern Azolla in 47 Ma Eocene sediments in an ACEX
Integrated Ocean Drilling Program (IODP) core in the middle of the Arctic
Ocean (Brinkhuis et al., 2006) adds an almost surreal element to this
vignette of crocodile-infested subtropical swamp forests on the shores of a
warm, fresh arctic ocean covered with floating aquatic plants. Studies of
such ice-free, high-latitude, deep-time analogues are important scientific
windows into how the Arctic ecosystem might operate in the absence of
permanent sea ice or in fully deglaciated conditions.
In a world with forested poles and tropical midlatitudes, the nature of the
equatorial realm is a serious question. Recent estimates of sea surface tem-
peratures for the Late Cretaceous to Eocene tropics, based on well-preserved
marine microfossils, suggest that temperatures may have exceeded 35-40ºC
(Huber, 2002; Norris et al., 2002; Pearson et al., 2007)—the absence of
equatorial coral reefs may have been because seawater was too hot. Ample
evidence from equator to pole shows that the last greenhouse was a very dif-
ferent place from today, and that the composition and distribution of biomes
were wholly different from the present—it was not just a warmer world, but
rather a completely different world from the present day.

OCR for page 26
62 UNDERSTANDING EARTH’S DEEP PAST
be at least two orders of magnitude higher than the long-term average
(Hassan et al., 2005), a rate potentially commensurate with the largest
mass extinctions of the geological past (Sepkoski, 1996; Bambach, 2006).
Modeling future biodiversity losses and their effects on the Earth’s eco-
systems and climate, however, is inherently difficult (Botkin et al., 2007),
making it imperative to assess the outcome of equivalent “natural experi -
ments” in the geological record (NRC, 1995; Myers and Knoll, 2001). The
five major, and dozens of minor, mass extinctions of the past half-billion
years (Sepkoski, 1996; Bambach, 2006) offer unique insights regarding eco-
system susceptibility and response to environmental stress, the potential
for ecological collapse, and the mechanisms of ecosystem recovery (Benton
and Twitchett, 2003; Bottjer et al., 2008). Furthermore, the integration of
paleontologic, stratigraphic, and geochemical records for many intervals
of the past half-billion years have revealed the variable character of past
biotic turnovers and mass extinction events (e.g., Boxes 2.4, 2.6, 2.7, 2.8),
which differ in regard not only to severity but also to duration, selectivity,
and the nature of environmental stresses (e.g., the transition out of super-
greenhouse conditions into Ordovician glaciation [Trotter et al., 2008]; the
Early to Middle Triassic radiations [Payne et al., 2004]; the nannoplankton
crisis and foraminiferal turnovers of the Cretaceous ocean anoxic events
[Leckie et al., 2002]; Eocene-Oligocene faunal extinction and immigration
[Kobashi et al., 2001; Ivany et al., 2004]). Most importantly, the geologi -
cal record uniquely captures past climate-ecological interactions that are
fully played out and thereby archive the impact, response, interaction, and
recovery from past global warming and major climate transitions.